A Dynamic Marine Calcium Cycle During the Past 28
Million Years
Elizabeth M. Griffith, et al.
Science 322, 1671 (2008);
DOI: 10.1126/science.1163614
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This intramolecular looping mechanism resembles that of a ring-closure reaction (MP→Ca*
in Fig. 3E), and the threading rate will therefore
depend on the effective molarity (EM) of the
reactive components (the cavity and the open
chain end). The EM (23) describes the relative
ease of ring-closure reactions compared with intermolecular reactions under otherwise identical conditions and has been tabulated in the literature (24).
EM values contain all the intrinsic parameters of
chains relevant for ring closure such as ring strain
and torsional entropy. In line with intuition, EM
values become smaller when the chains connecting
the reactive components become longer. According to the model (Fig. 3E), the intramolecular path
will be more favored than the intermolecular path
as long as KaEM > 1, in which Ka is the association
constant of the outside-viologen complex. At the
given values of Ka (≈40 M−1), these conditions
are met for chains shorter than 40 atoms (because
EM = 0.025 M for chains of 40 atoms), which is
in line with the observation that the polymer
chains (90 to 440 atoms) apparently do not benefit from this looping process. The intramolecular
process will not only be lower in transition state
energy than the intermolecular process but also
energetically downhill in the direction of the viologen trap. In the process of moving along the
chain toward the viologen moiety (Ca* → Cb* →
Cc* in Fig. 3E), the chain length between the
reactive components (the outside of the macrocycle
and the viologen moiety) decreases, and the EM
value consequently increases (EMa < EMb < EMc),
resulting in a downhill reaction coordinate (Fig.
3C). The resulting unidirectional movement continues until the loop becomes too small (Cc*, Fig. 3,
C and E) and the viologen dissociates from the
outside of the macrocycle, after which the last eight
atoms are traversed in a nonstabilized fashion and
the macrocycle embraces the viologen trap. As a
result, the traversing of the last eight atoms becomes
rate-limiting, which is in agreement with the
experimental results that the observed rates never
become significantly higher than the rate of
traversing a chain of eight atoms (Fig. 2).
To further support the model, we synthesized two
different porphyrin macrocycles (2-H2 and 2-Zn) (25)
with -OH functionalities on the outside of the cavity, which form stronger outside complexes with
Vdb [Ka = 2 × 104 M−1 and 5 × 104 M−1 for 2-Zn and
2-H2, respectively; see figs. S9, S10, and accompanying text (16)]. This modification was expected
to result in higher threading rates over the polymer
chains as well because KaEM > 1 in a larger region
of chain lengths. It should be noted that, even at
these high association constants, quenching caused
by the outside complexes is less than 5% of the total
measured quenching at the applied experimental
concentrations (fig. S11). The fluorescence threading studies of 2-H2 and 2-Zn revealed that both 2-H2
and 2-Zn traverse the short chains and the oligomer
chains in an analogous fashion to 1-H2 and 1-Zn
(Fig. 2 inset and Table 1). However, the threading
over the polymer chains is significantly faster for
both 2-H2 (on average 2.2 times) and 2-Zn (up to a
factor of 4). These results thus support the idea that
the higher affinity of 2-H2 and 2-Zn for the viologen
moiety on the outside of the macrocycle enhances
the rates of threading over the polymer chains via
the intramolecular looping process.
The experimentally determined length dependencies for the threading process can be described
by a modified consecutive hopping model, which
includes a kinetically favorable so-called entron
effect. The above results highlight that weak supramolecular interactions between the outside of the
macrocycle and the chain have a dramatic effect
on the overall threading process. It is a classical
example of how small interactions can lead to accelerated rates by means of transition state stabilization. In our simple biomimetic model system,
these interactions both initiate and guide the threading process, resulting in accelerated rates and
mimicking the natural translocation systems in
which interactions between the biopolymer and a
receptor in the vicinity of the pore are fundamental
and essential steps in efficient and selective translocation (12, 13). Given that our studies show that
even weak interactions between the macrocyclic
exterior and the polymer chains already cause
significant increases in the threading rates, it is
very likely that such assisted threading mechanisms not only play a role in the specific natural
examples but also in a much larger number of
natural polymer translocation systems. Another
intriguing effect of the mechanism is that unidirectional motion (achieved without the application of an external driving force) is facilitated
toward the trapping site. This driving force should
allow for the design of novel artificial machines
(26) and catalysts in which the threading can be
finely controlled.
References and Notes
1. W. Wickner, R. Schekman, Science 310, 1452 (2005).
2. E. J. Mancini et al., Cell 118, 743 (2004).
3. A. Johnson, M. O’Donnell, Annu. Rev. Biochem. 74, 283
(2005).
4. E. R. Kay, D. A. Leigh, F. Zerbetto, Angew. Chem. Int. Ed.
46, 72 (2007).
5. R. A. Bissell, E. Cordova, A. E. Kaifer, J. F. Stoddart,
Nature 369, 133 (1994).
6. A. M. Brouwer et al., Science 291, 2124 (2001); published
online 22 February 2001 (10.1126/science.1057886).
7. J. D. Badjic, V. Balzani, A. Credi, S. Silvi, J. F. Stoddart,
Science 303, 1845 (2004).
8. E. R. Kay, D. A. Leigh, Nature 440, 286 (2006).
9. M. C. Jiménez, C. Dietrich-Buchecker, J.-P. Sauvage,
Angew. Chem. Int. Ed. 39, 3284 (2000).
10. Y. Liu et al., J. Am. Chem. Soc. 127, 9745 (2005).
11. P. Thordarson, E. J. A. Bijsterveld, A. E. Rowan, R. J. M. Nolte,
Nature 424, 915 (2003).
12. G. Blobel, ChemBioChem 1, 86 (2000).
13. T. A. Rapoport, Nature 450, 663 (2007).
14. R. G. E. Coumans, J. A. A. W. Elemans, R. J. M. Nolte,
A. E. Rowan, Proc. Natl. Acad. Sci. U.S.A. 103, 19647
(2006).
15. P. Hidalgo Ramos et al., J. Am. Chem. Soc. 129, 5699 (2007).
16. Materials and methods are detailed in supporting online
material available at Science Online.
17. W. Herrmann, B. Keller, G. Wenz, Macromolecules 30,
4966 (1997).
18. A. Harada, J. Li, M. Kamachi, Nature 356, 325 (1992).
19. L. P. Hammett, Physical Organic Chemistry (McGraw-Hill,
New York, ed. 2, 1970), p. 112.
20. A. B. Kolomeisky, Biophys. J. 94, 1547 (2008).
21. E. Slonkina, A. B. Kolomeisky, J. Chem. Phys. 118, 7112
(2003).
22. A. Meller, L. Nivon, D. Branton, Phys. Rev. Lett. 86, 3435
(2001).
23. L. Mandolini, Adv. Phys. Org. Chem. 22, 1 (1986).
24. C. Galli, L. Mandolini, Eur. J. Org. Chem. 3117 (2000).
25. P. Thordarson et al., J. Am. Chem. Soc. 125, 1186 (2003).
26. W. R. Browne, B. L. Feringa, Nat. Nanotechnol. 1, 25 (2006).
27. This research was supported by a National Research
School Combination Catalysis Controlled by Chemical
Design grant, Nederlandse Organisatie voor
Wetenschappelijk Onderzoek Veni, Vidi and Vici grants
(J.A.A.W.E. and A.E.R.), a Koninklijke Nederlandse
Akademie van Wetenschappen grant (R.J.M.N.), and a
Nanoned grant (A.E.R. and R.J.M.N.).
Supporting Online Material
www.sciencemag.org/cgi/content/full/322/5908/1668/DC1
Materials and Methods
Figs. S1 to S18
Table S1
References
14 August 2008; accepted 14 November 2008
10.1126/science.1164647
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REPORTS
A Dynamic Marine Calcium Cycle
During the Past 28 Million Years
Elizabeth M. Griffith,1,2* Adina Paytan,2 Ken Caldeira,3 Thomas D. Bullen,4 Ellen Thomas5,6
Multiple lines of evidence have shown that the isotopic composition and concentration of
calcium in seawater have changed over the past 28 million years. A high-resolution, continuous
seawater calcium isotope ratio curve from marine (pelagic) barite reveals distinct features in the
evolution of the seawater calcium isotopic ratio suggesting changes in seawater calcium
concentrations. The most pronounced increase in the d44/40Ca value of seawater (of 0.3 per mil)
occurred over roughly 4 million years following a period of low values around 13 million years
ago. The major change in marine calcium corresponds to a climatic transition and global change in
the carbon cycle and suggests a reorganization of the global biogeochemical system.
he marine calcium (Ca2+) cycle is related
to processes that control oceanic alkalinity and atmospheric CO2, playing an important role in earth’s climate (1). Calcium
carbonate (CaCO3) sedimentation in the ocean
T
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VOL 322
represents the largest carbon sink in the combined atmosphere, biosphere, and ocean system, and thus strongly influences the global
carbon cycle (2). Fluctuations in CaCO3 sedimentation reflect imbalances in the long-term
12 DECEMBER 2008
1671
carbon cycle and are related to global changes
in climate and tectonics (3). Such fluctuations
may also be reflected in the marine Ca2+ cycle
and seawater Ca2+ concentrations and isotopic
ratios (4, 5).
Changes in Ca2+ concentration of seawater
throughout the Phanerozoic are revealed in
brine inclusions (6, 7), elemental ratios in marine skeletal carbonates (8), and the changing
abundance of carbonate and potash-evaporite
minerals (9, 10). Shifts in seawater Ca2+ concentration during the past 28 million years (My) are
thought to reflect predominantly sedimentological factors (e.g., carbonate deposition and dolomitization) (11), although changes in sea-floor
generation rates may also play a role (12).
The amount of Ca2+ in the modern ocean and
its isotopic composition [d44/40Caseawater (13)]
are determined by the balance between riverine and hydrothermal inputs and removal through
CaCO3 deposition and alteration of oceanic crust
and their respective isotopic compositions (fig. S1).
Estimates for the isotopic composition of Ca2+ inputs from continental rivers range from d44/40Ca =
–0.87 to –1.30 per mil (‰) (5, 14–16); hydrothermal inputs average –0.96 T 0.20‰ (15).
When CaCO3 is precipitated from seawater, its
Ca-isotopic composition is 0.7 to 1.6‰ less than
that of dissolved Ca2+ in seawater, depending
on which mineral, calcite or aragonite, is precipitated and on parameters such as temperature and
precipitation rate (17). Alteration of oceanic crust,
estimated to account for less than 10% of the Ca2+
sink in the modern ocean, has an average d44/40Ca
between 0.7 and 1.7‰ less than d44/40Caseawater
(18). Due to the range of isotopic signatures in the
modern ocean and the uncertainty of the representative average values, it is not known precisely
whether or over what time scale the marine Ca2+
cycle is in isotopic steady state (5, 14, 15, 19).
Recent work using Ca-isotopic ratios measured
in marine carbonates (14, 16, 19–21) has shown
that the isotopic signature recorded in carbonates
fluctuated over time. However, there is no consensus as to whether seawater composition or environmental and physiological factors ultimately
control the Ca-isotopic composition measured in
marine carbonates (16, 19). We present a highresolution, continuous seawater Ca-isotope curve
from marine (pelagic) barite (BaSO4) to help ex-
plain factors controlling the cycles of calcium and
carbon over the past 28 My.
Marine barite, a minor component of marine
sediments, is a useful recorder of changes in seawater chemistry (22–24). Calcium (Ca2+) substitutes for barium (Ba2+) in the barite crystal lattice,
providing an archive of Ca-isotopes in a highly
stable sulfate mineral (25). Marine barite has advantages over carbonate minerals for studying the
seawater Ca-isotopic ratio because of its resistance
to diagenesis in oxic pelagic sediments and its uninterrupted record over important climate intervals
associated with carbonate dissolution (22). In addition, marine barite precipitates inorganically in
seawater; thus, no physiological “vital effects” are
expected (22).
The measured d44/40Ca of Holocene marine
barite separated from more than 20 core-top deep
sea sediment samples in the major ocean basins are
indistinguishable from each other, with a d44/40Ca
of –2.01 T 0.15‰ (average 2smean) (25). The offset from seawater (isotopic fractionation) is not
strongly related to any measured environmental
parameter (25) and encompasses the expected range
of conditions over the past 28 My at the location of
our down-core samples. This offset can therefore
be considered constant during the past 28 My (26).
The seawater Ca-isotope curve for the past
28 My was reconstructed from marine barite
at a resolution better than 1 My from 51 sediment
samples collected by the Deep Sea Drilling
Project (DSDP) and the Ocean Drilling Program
(ODP) at sites 572 to 575, 1218, and 1219 in
the east equatorial Pacific (27).
The marine barite d44/40Caseawater record
(Fig. 1) displays a long-term trend toward more
positive values, in general agreement with existing seawater Ca-isotope records from biogenic carbonates (Fig. 1 and table S1). However,
our data also reveal features that correspond
with shifts in the marine carbon and oxygen
isotope records (Fig. 1). The d44/40Caseawater remained fairly stable at around –0.3‰ T 0.1‰
(2smean, n = 26) from 28 million years ago (Ma)
until around 13 Ma. From 13 Ma to 8 Ma, it
rose to the modern value, where it remained
until the present (0 T 0.1‰, 2smean, n = 23).
We use a numerical model to investigate
the implications of this Ca-isotope record for
the marine calcium biogeochemical cycle in
relation to major climatic events and perturbations in the carbon cycle [following (14);
for details, see (27)]. The state variables in the
model are the isotopic composition of seawater Ca2+ (model input from the barite data)
and its concentration. Model parameters include the isotopic signature of the combined
input flux of Ca2+ (riverine, hydrothermal, and
dolomitization) and the isotopic fractionation
during CaCO3 precipitation. The oceanic residence
time of Ca2+ is taken to be 1.3 My (21) and
assumed to be constant (28). Using a constant
sedimentation flux (and changing residence
time) did not drastically vary the pattern in calculated Ca2+ concentrations (27).
An initial condition for 28 Ma is constructed
in the model, and the set of equations are solved
simultaneously, so the system is allowed to evolve
Fig. 1. Ca-isotopic composition of
seawater (d44/40Caseawater) over the
past 28 My. All the Ca-isotopic
ratios are reported in per mil relative to seawater and are plotted
with respect to the chronology in
(33) [see (27)]. Error bars are the
precision of each sample calculated
as the 2smean of replicate analyses
on the thermal ionization mass
spectrometer or the average 2smean
if only one analysis was made. The
bold solid black curve is the cubic
smoothing spline of the data (34)
with T 0.18‰ (average 2smean) in
gray. Two samples with 2smean >
0.4‰ were removed from the trend
and figure. Averages of duplicate
samples are plotted. Smoothed
global d18O and d13C synthesis of
benthic (deep sea) foraminifera reported relative to Pee Dee belemnite
in per mil (35). Pleistocene, Ple.;
Pliocene, Plio.
1
Department of Geological and Environmental Sciences,
Stanford University, Building 320, Room 118, Stanford, CA
94305, USA. 2Institute of Marine Sciences, University of
California, Santa Cruz, CA 95064, USA. 3Department of
Global Ecology, Carnegie Institution of Washington, 260
Panama Street, Stanford, CA 94305, USA. 4Branch of Regional Research, Water Resources Division, U.S. Geological
Survey, MS 420, 345 Middlefield Road, Menlo Park, CA
94025, USA. 5Department of Geology and Geophysics, Yale
University, Post Office Box 208109, New Haven, CT 06520,
USA. 6Department of Earth and Environmental Sciences,
Wesleyan University, 265 Church Street, Middletown, CT
06459, USA.
*To whom correspondence should be addressed. E-mail:
egriffith@ucsc.edu
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REPORTS
toward the present. In the model, the assumption of
constant average values estimated for the presentday Ca2+ system (sources and sinks) (fig. S1) leads
to unacceptable seawater Ca2+ fluctuations. Accordingly, the isotopic composition of the input
flux of Ca2+ (d44/40Cain) and/or the isotopic fractionation associated with the sink (D44/40Cased) must
have varied over this time period. First, we varied
d44/40Cain with time (Fig. 2A), constrained at points
where the Ca2+ concentration is known from measurements of brine inclusions (7) and linearly extrapolated between these points as described in (27).
Because the first approach assumes timedependent isotopic variation in the input flux
(Fig. 2A), the general long-term trend in the Caisotopic composition of seawater is directly related to changes in the isotopic signature of this
flux. With these assumptions, a shift in d44/40Cain
to more positive values likely reflects a shift
in the weathered terrain to more continental
sourced silicate rocks with a more positive
d44/40Ca value (16). An increase from 6% to
17% in the contribution of silicate weathering
to the total weathering flux can explain the
data (27). The d44/40Ca value of carbonates subjected to weathering also changes, contributing
about 0.1‰ of the 0.6‰ long-term increase
in the modeled d44/40Cain over this time period (20).
Dolomitization (a source of Ca2+ to the oceans)
could have influenced d44/40Cain in the past.
Dolomitization releases Ca2+ by diffusional ex-
change with Mg2+, adding Ca2+ depleted in
Ca without affecting the seawater carbonate
flux (19). The rate of marine dolomite formation has decreased over the Cenozoic to a present
minimum (11), consistent with the increasing
d44/40Ca seawater over the past 28 My, but it is
unclear how important this process has been
as a source of Ca2+ during this time interval or
over the observed transition about 13 Ma.
In addition, the isotopic fractionation associated with the sink (D44/40Cased) likely changed.
Shifts in the D44/40Cased could have been caused
by changes in temperature during precipitation, rate of precipitation, or in the mineralogy
of the deposited carbonates (aragonite versus
calcite), which is controlled by the dominant
calcifying organisms (10, 20). In the second approach (Fig. 2B), D44/40Cased varies and is calculated as the difference between seawater
values in this study and the bulk carbonate
(output) record of (21). d44/40Cain also varies,
as in the first model, and is likewise constrained
by the long-term trend in seawater Ca2+ concentration from fluid inclusions (27). This results
in a slightly different d44/40Cain reconstruction
due to the differences in D44/40Cased histories
we used. Changing the D44/40Cased in addition
to d44/40Cain does not in general change the
shape of the modeled seawater Ca2+ concentration curve.
In summary, model results indicate that the
isotopic composition of the combined input
44
Fig. 2. Time varying isotopic composition of the calcium input flux to the oceans (d44/40Cain) and
isotopic fractionation during calcium carbonate precipitation before sedimentation (D44/40Cased) in
per mil used in the numerical models (top). Calculated ratio of amount of calcium in the oceans
[NCa(t)], relative to the present value [NCa(modern)] constrained using data (open diamonds) from
(7). Solid black line is NCa(t) / NCa(modern); dashed black line is Phanerozoic trend (7) (bottom).
Numerical model assuming (A) constant D44/40Cased or (B) varying D44/40Cased calculated from the
difference between seawater d44/40Ca (this study) and measured contemporaneous bulk nannofossil
ooze carbonates (14, 21).
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flux of Ca2+ and/or isotopic fractionation during carbonate precipitation must have changed
over the past 28 My (21, 16), and these results
suggest changes in seawater Ca2+ concentrations. Specifically, the relatively abrupt (<5 My)
changes in the d44/40Caseawater curve indicate
large, rapid transient changes in the amount of
Ca2+ in the ocean, of up to twice the present
value (Fig. 2), and/or transient fluctuations in
the Ca-isotopic composition of the input or sedimentation fluxes of up to 1‰. However, such
large and abrupt changes in the Ca-isotopic
compositions of these input or output fluxes is
outside the range of measured present-day values (5, 14–16) and does not seem to be a likely
mechanism for explaining all variation in our
seawater Ca-isotope record.
The most prominent change in the oceanic
Ca2+ concentration corresponds to a major climatic transition at ~13 Ma (middle Miocene),
when the earth’s climate changed from a period of relative global warmth into a cooler
climate, while the Antarctic ice sheet increased
in volume (29). This increase in Ca2+ concentration may be a consequence of an increase in
the Ca2+ flux to the ocean.
An increase in the Ca2+ flux from continental weathering (without an accompanying
increase in carbonate deposition) from recently
exposed continental shelves, increased weathering, and/or transport of weathered material into
the oceans associated with glaciation might have
resulted from a drop in sea level as shown by
the sharp 0.5‰ increase in the d18O of benthic
foraminifera (30) (Fig. 1). Increased weathering
of carbonate sediments (relative to silicate sediments) increases the flux of Ca2+ to the ocean
relative to the carbon (C) flux providing a potential means for decoupling the Ca2+ and C systems (19). The rate of change in the Ca2+ flux
must be decoupled from and higher than the rate
of change in the flux of C; otherwise, Ca2+ flux
variations are quickly (<<1 My) buffered by feedbacks involving terrestrial rock weathering and
marine C-CaCO3 deposition/dissolution processes.
The calcium carbonate compensation depth
(CCD), the water depth where the calcite rain
rate is balanced by its dissolution rate such
that all calcite produced in the water column is
dissolved, records the balance between CaCO3
deposition/dissolution in deep sea sediments
(31). Shifts in the CCD occurred during the past
28 My (31), resulting from changes in deep
ocean chemistry and the supply rate of CaCO3
to the sediments, but long-term changes in the
equatorial Pacific CCD (>1 My) do not appear
to be solely controlled by oceanic Ca2+ concentrations (fig. S7).
A change in the marine carbon cycle occurred
during the middle Miocene climatic transition,
when the carbon isotopic composition of bulk
carbonates (d13Ccarb) began to decrease by about
2‰ to modern values (30) (Fig. 1). If the total
C flux to the ocean had been constant, this would
have implied a shift to more continental-sourced
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REPORTS
1673
rocks rich in organic C relative to carbonate
weathering over the past ~13 My (31). Previous interpretations of 87Sr/86Sr, 187Os/186Os,
and d7Li records agree with a change in the
nature of weathered material at this time (32).
Many aspects of the biogeochemical Ca2+
cycle need to be constrained better by observational and experimental data and paleoceanographic and geologic evidence before a unique
interpretation of the causes of the long-term
trend and abrupt transitions in d44/40Caseawater is
possible. Reconstructions of the global CCD
and high-resolution records of seawater Ca2+
concentrations (e.g., from fluid inclusions or
other proxies) and of the rate of dolomitization
and sea-floor generation are all needed.
Our identification of an abrupt middle Miocene change in the d44/40Caseawater record from
marine barite attests to the dynamic nature of an
element deemed “conservative” in the present
ocean and demonstrates that we cannot assume
that concentrations of Ca2+ in seawater have
been stable over the Cenozoic.
References and Notes
1. H. C. Urey, The Planets, Their Origins and Development
(Yale Univ. Press, New Haven, CT, 1952).
2. A. J. Ridgwell, M. J. Kennedy, K. Caldeira, Science 302,
859 (2003).
3. M. I. Budyko, A. B. Ronov, A. L. Yanshin, History of the
Earth's Atmosphere (Springer-Verlag, New York, 1987).
4. J. Skulan, D. J. DePaolo, T. L. Owens, Geochim.
Cosmochim. Acta 61, 2505 (1997).
5. P. Zhu, J. D. Macdougall, Geochim. Cosmochim. Acta 62,
1691 (1998).
6. T. K. Lowenstein, M. N. Timofeeff, S. T. Brennan,
L. A. Hardie, R. V. Demicco, Science 294, 1086 (2001).
7. J. Horita, H. Zimmermann, H. D. Holland, Geochim.
Cosmochim. Acta 66, 3733 (2002).
8. J. A. D. Dickson, Science 298, 1222 (2002).
9. P. A. Sandberg, Nature 305, 19 (1983).
10. L. A. Hardie, Geology 24, 279 (1996).
11. H. D. Holland, Am. J. Sci. 305, 220 (2005).
12. C. P. Conrad, C. Lithgow-Bertelloni, Geology 35, 29 (2007).
13. The Ca-isotopic composition of a sample is expressed as
the deviation from a standard solution value, using
d-notation in per mil: d44/40Ca = [(44Ca/40Ca sample –
44
Ca/40Ca standard) / 44Ca/40Ca standard] × 1000, where
44
Ca/40Ca standard refers to the 44Ca/40Ca isotopic
ratio of seawater.
14. C. L. De La Rocha, D. J. DePaolo, Science 289, 1176 (2000).
15. A. D. Schmitt, F. Chabaux, P. Stille, Earth Planet.
Sci. Lett. 213, 503 (2003).
16. N. G. Sime et al., Geochim. Cosmochim. Acta 71, 3979
(2007).
17. N. Gussone et al., Geochim. Cosmochim. Acta 69, 4485
(2005).
18. M. Amini et al., Geochim. Cosmochim. Acta 72, 4107 (2008).
19. A. Heuser et al., Paleoceanography 20, 10.1029/
2004PA001048 (2005).
20. M. S. Fantle, D. J. DePaolo, Earth Planet. Sci. Lett. 237,
102 (2005).
21. J. Farkaš et al., Geochim. Cosmochim. Acta 71, 5117 (2007).
22. A. Paytan, M. Kastner, E. E. Martin, J. D. Macdougall,
T. Herbert, Nature 366, 445 (1993).
23. A. Paytan, M. Kastner, D. Campbell, M. H. Thiemens,
Science 304, 1663 (2004).
24. A. V. Turchyn, D. P. Schrag, Science 303, 2004 (2004).
25. E. M. Griffith, E. A. Schauble, T. D. Bullen, A. Paytan,
Geochim. Cosmochim. Acta 72, 5641 (2008).
26. Range of parameters measured include conditions
relevant to the ranges for the measured samples
(DSDP-ODP sites): Holocene core-top sample sites include
the average annual temperature in the upper 700 m
of the water column, where the majority of marine barite
is thought to precipitate, 1 to 14°C; the depth of each
core-top location, 3175 to 4431 m; water column
barite saturation at sea floor, SI = 0.6 to 0.8 and at
700 m, SI = 0.8 to 1.2; average upper (700 m) water
column: salinity (34.2 to 34.9), carbonate concentration
(86.65 to 139.56 mmol/kg), and dissolved oxygen
concentration (1.1 to 6.2 ml/l); bulk sedimentation
rate (1.3 to 2.7 cm/thousand years) and barite accumulation
rates (0.7 to 23.4 mg m–2 year–1).
27. Materials and methods are available as supporting
material on Science Online.
28. A constant residence time in the model requires that
the output flux of Ca2+ and the total amount of dissolved
Ca2+ in the oceans vary together. This is a valid
assumption if changes in Ca2+ concentration occurred
due to changes in the input flux of Ca2+ to the ocean that
Earthquake Supercycles Inferred from
Sea-Level Changes Recorded in the
Corals of West Sumatra
Kerry Sieh,1* Danny H. Natawidjaja,2 Aron J. Meltzner,1 Chuan-Chou Shen,3
Hai Cheng,4 Kuei-Shu Li,3 Bambang W. Suwargadi,2 John Galetzka,1
Belle Philibosian,1 R. Lawrence Edwards4
Records of relative sea-level change extracted from corals of the Mentawai islands, Sumatra,
imply that this 700-kilometer-long section of the Sunda megathrust has generated broadly
similar sequences of great earthquakes about every two centuries for at least the past 700 years.
The moment magnitude 8.4 earthquake of September 2007 represents the first in a series of
large partial failures of the Mentawai section that will probably be completed within the next
several decades.
arge sections of the great arcuate fault beneath the eastern flank of the Indian Ocean
have failed progressively over the past 8
years in an extraordinary sequence of big earthquakes (Fig. 1) (1–4). The largest of these failures
L
1674
of the Sunda megathrust, in 2004, caused the most
devastating tsunami the world has seen in many
generations. One question of great humanitarian
and scientific importance is which remaining
unruptured sections of the megathrust will fail next.
12 DECEMBER 2008
VOL 322
SCIENCE
29.
30.
31.
32.
33.
34.
35.
36.
are not accompanied with changes in the C cycle. If a
change in Ca2+ input flux was associated with and
compensated by a change in CaCO3 sedimentation
through an increase in alkalinity, no change in Ca2+
concentration would result. Increasing or decreasing the
residence time by 50% gives similar fluctuations but
different magnitudes (peak concentrations range from
180% to 290% of present values) (fig. S5). Modeling the
Ca2+ system with full carbonate chemistry reveals the
sensitivity of the Ca2+ cycle to changes in C on long time
scales (>10,000 years) (27). A mechanism for decoupling
the two cycles is required in order to result in changes in
Ca2+ concentration. These might include changes in
the rate of dolomitization or the Ca:bicarbonate ratio in
the riverine flux due to shifts in the weathering regime.
A. Holbourn, W. Kuhnt, M. Schulz, H. Erlenkeuser, Nature
438, 483 (2005).
M. Lyle, Paleoceanography 18, 10.1029/2002PA000777
(2003).
D. A. Hodell, P. A. Mueller, J. A. McKenzie, G. A. Mead,
Earth Planet. Sci. Lett. 92, 165 (1989).
E. C. Hathorne, R. H. James, Earth Planet. Sci. Lett. 246,
393 (2006).
L. Lourens, F. Hilgen, N. J. Shackleton, J. Laskar, D. Wilson,
in A Geological Time Scale 2004, F. Gradstein, J. Ogg,
A. G. Smith, Eds. (Cambridge Univ. Press, Cambridge, UK,
2004), pp. 409–440.
Generated using a smoothing parameter (P = 0.80) and
data weighted by its precision in the MATLAB 7.6.0.324
(R2008a) cubic smoothing spline function csaps.
J. Zachos, M. Pagani, L. Sloan, E. Thomas, K. Billups,
Science 292, 686 (2001).
Samples provided by the Integrated Ocean Drilling
Program. We thank A. Eisenhauer, J. Fitzpatrick,
W. A. Griffith, R. Jones, and G. Li for analytical assistance.
Supported by NSF CAREER Grant OCE-0449732 (A.P.)
and by a National Defense Science and Engineering
Graduate Fellowship and an NSF Graduate Research
Fellowship (E.M.G.).
Supporting Online Material
www.sciencemag.org/cgi/content/full/322/5908/1671/DC1
Materials and Methods
Figs. S1 to S7
Table S1
References
22 July 2008; accepted 7 November 2008
10.1126/science.1163614
Until late 2007, the largest remaining unbroken Sumatran section had been the 700-kmlong Mentawai patch, dormant since two great
earthquakes in 1797 and 1833 (5). Modeling of
coral and instrumental geodetic data had shown
most of the patch to have been highly coupled
throughout at least the past half-century (6, 7).
That is, overlying and underlying blocks have
been locked together, and strains continue to
accumulate. These observations had led to concerns that the remainder of the Mentawai patch
might rupture soon (8–11). In September 2007,
rapid-fire failure of portions of the Mentawai
patch produced a moment magnitude (Mw) 8.4
earthquake and several large aftershocks (12),
which served to further elevate anxiety.
1
Tectonics Observatory, California Institute of Technology,
Pasadena, CA 91125, USA. 2Research Center for Geotechnology, Indonesian Institute of Sciences, Bandung, Indonesia.
3
Department of Geosciences, National Taiwan University,
Taipei, Taiwan, ROC. 4Department of Geology and Geophysics,
University of Minnesota, Minneapolis, MN 55455, USA.
*To whom correspondence should be addressed at the
Earth Observatory of Singapore, Nanyang Technological
University, 639798 Singapore. E-mail: sieh@ntu.edu.sg
www.sciencemag.org
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